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Retrieving timescales of oceanic crustal evolution at
Oceanic Core Complexes: Insights from diffusion
modelling of geochemical profiles in olivine
Carlotta Ferrando, Kendra Lynn, Valentin Basch, Benoit Ildefonse,
Marguerite Godard
To cite this version:
Carlotta Ferrando, Kendra Lynn, Valentin Basch, Benoit Ildefonse, Marguerite Godard.
Re-trieving timescales of oceanic crustal evolution at Oceanic Core Complexes: Insights from
dif-fusion modelling of geochemical profiles in olivine.
Lithos, Elsevier, 2020, 376-377, pp.105727.
Retrieving
timescales of oceanic crustal evolution at Oceanic Core
Complexes:
Insights from diffusion modelling of geochemical
pro
files in olivine
Carlotta
Ferrando
a,b,
Kendra
J. Lynn
c,
Valentin Basch
b,
Benoit Ildefonse
a,
Marguerite Godard
aaGéosciences Montpellier, Université de Montpellier, CNRS, Université des Antilles, Montpellier, France b Dipartimento di Scienze della Terra e dell'Ambiente, University of Pavia, Pavia, Italy
c Department of Earth Sciences, University of Delaware, Newark, DE 19713, USA
a b s t r a c t
Keywords:
Lower oceanic crust at the Atlantis Massif Olivine trace elements
Diffusive re-equilibration Oceanic crustal cooling Magma emplacement
Exhumation by detachment faults
The building of oceanic crust at Oceanic Core Complexes (OCC) has been described as a complex process involv-ing multiple intrusions of magma over a protracted period of time. The migration of primitive magmas (i.e., Mid Ocean Ridge Basalts, MORBs) can lead to melt-rock interactions during reactive porousflow processes through the lithosphere. The timescales of these reactive processes and the subsequent cooling of the modified crystal matrix remain unconstrained. Diffusion modelling has been widely used to retrieve timescales of magmatic pro-cesses. In this study, we use diffusion models to constrain (i) the minimum timescales of melt-rock interactions and (ii) the cooling rates of the gabbroic sequence forming the oceanic crust at an OCC. We chose samples of the most primitive olivine-rich troctolites from the gabbroic sequence sampled in IODP Hole U1309D (Atlantis Massif OCC, Mid-Atlantic Ridge 30°N). Olivine-rich troctolites were interpreted as marking local partial assimilation of mantle intervals into the oceanic crust, and thus allowed to understand the dynamics of mantle assimilation and the formation of slow-spreading oceanic crust at the Atlantis Massif OCC. Olivines in olivine-rich troctolites represent relicts of pre-existing mantle olivine, while clinopyroxenes and plagioclases are crystallized during reactive percolation. Olivine chemical compositions show that olivine-rich troctolites inherit chemical heteroge-neity from the mantle precursor. Flat geochemical profiles in olivine indicate complete chemical re-equilibration of olivine crystals with the locally modified percolating melt. Exception is made for most Ca profiles that show lower Ca contents at the olivine rim compared to the relative crystal core, as the result of subsolidus cooling. Three-dimensional (3D) diffusion models at magmatic conditions (T = 1210–1300 °C and P = 2 kbar) reveal that complete chemical re-equilibration of 4 mm-size mantle-derived olivine with percolating MORB-type melts can be attained within durations of less than 300 yr. The Ca-in-olivine geospeedometer reveal that cooling rates from ~1200 °C to ~1050 °C are constant downhole and on average 0.004 °C/yr; they are comparable with lower temperature cooling rates (850 °C–250 °C) estimated at the Atlantis Massif OCC. The minimum timescales from 3D models point to rather fast re-equilibration of olivine. The downhole chemical heterogeneity inherited from the precursor mantle, coupled with the timescales of diffusive re-equilibration suggest that the partial as-similation of the upwelling mantle and its incorporation into the oceanic crust occurred in the time-frame of a single melt input. Our cooling data are consistent with building of the oceanic crust at OCCs controlled by contin-uous uplift, in turn governed by long-lived detachment faults. The latter contribute to the rapid cooling of the as-similated mantle intervals and the magma bodies. Diffusion models of geochemical profiles in olivine from a single crustal section allow to reconstitute the early magmatic processes leading to mantle assimilation and early crystallization of gabbros, and the cooling history of the oceanic crust at OCC from magmatic conditions to hydrothermalism.
1. Introduction
Melts generated by adiabatic melting of the upwelling mantle be-neath mid-ocean ridges are supplied to the magmatic plumbing system, where they undergo complex processes of magma crystallization, mixing and/or assimilation of a pre-existing crystal matrix prior to
eruption on the seafloor (e.g.,Brandl et al., 2016;Coogan and O'Hara, 2015;O'Hara, 1977;O'Neill and Jenner, 2012). These processes are re-corded in the lower oceanic crust beneath oceanic spreading centers. Recent studies have demonstrated the importance of reactive porous flow and melt-rock interactions in the construction of the gabbroic
lower crust, in particular along slow-spreading ridges (e.g.,Dick et al.,
2010;Drouin et al., 2009, 2010;Lissenberg and Dick, 2008;Lissenberg and MacLeod, 2016;Lissenberg et al., 2019;Sanfilippo et al., 2015)
and in their ophiolite analogues (e.g.,Basch et al., 2019a, 2019b, 2018;
Bédard, 1993, 2001;Bédard and Hébert, 1996;Rampone et al., 2020;
Sanfilippo et al., 2014). These reactive processes are thought to occur in cooling yet still partially molten magmatic systems, possibly at the transition between lithospheric mantle and the magmatic lower oceanic
crust (e.g.,Collier and Kelemen, 2010;Ferrando et al., 2018;Rampone
et al., 2020;Sanfilippo et al., 2015). To better understand the mecha-nisms driving the formation of slow-spreading oceanic crust, it is imper-ative to constrain the timescales of the processes involved in its evolution from early magmatic differentiation of melts (i.e., crystalliz-ation and dissolution-precipitcrystalliz-ation reactions) to cooling of the crustal sequence (i.e., crystal-dominated system).
Timescales of magmatic emplacement beneath slow-spreading ridges have been determined for crustal sections exposed along
the Mid-Atlantic Ridge (MAR; e.g.,Grimes et al., 2008), while duration
of cooling have been concurrently estimated by other authors for the lower oceanic crust beneath both the MAR and the Southwest
Indian Ridge (SWIR) (e.g., Grimes et al., 2011; John et al., 2004;
Schoolmeesters et al., 2012).John et al. (2004)documented the thermal structure of oceanic lithosphere beneath the ultraslow-spreading SWIR. They applied thermochronometric techniques on minerals in felsic veins (crystallization at ~850 °C), thus lacking direct constraints on the earlier magma emplacement history at higher temperatures (from ~1250 °C to ~850 °C). A broad reconstruction of timescales of oceanic crustal evolution from high temperature magma emplacement to cooling of the gabbroic sequence has not yet been established for a sin-gle suite of samples. In this contribution, we use the geochemical com-position of olivines from the widely investigated primitive olivine-rich
troctolites recovered at Atlantis Massif (IODP Hole U1309D;Blackman
et al., 2006) to provide timing of magmatic processes occurring at tem-peratures of about 1230 °C and of cooling of the lower oceanic crust down to about 1000 °C.
Olivine-rich troctolites from the oceanic crustal section at the
Atlan-tis Massif Oceanic Core Complex (OCC; IODP Hole U1309D;Blackman
et al., 2006) are one of the best studied olivine-rich lithologies (i.e., with >70% modal olivine) worldwide. They were interpreted to
result from partial mantle assimilation during reactive porousflow of
a MORB-type melt (Drouin et al., 2007, 2009, 2010;Ferrando et al.,
2018;Suhr et al., 2008). Olivine-rich troctolites from Hole U1309D were selected for this study because they are unique samples from ac-tive spreading ridges that record the oceanic crustal evolution from early magmatic processes of melt-rock interaction and crystallization of a primitive MORB, to exhumation and cooling during OCC formation. We apply diffusion modelling to provide a comprehensive dataset of timescales related to these magmatic processes and to reconstruct the geodynamic evolution of the oceanic crust from the Atlantis Massif OCC. Diffusion modelling of chemical species in rock-forming minerals is a tool widely used to quantify the timescales of magmatic processes
such as magma assimilation (e.g.,Bindeman et al., 2006;Costa and
Dungan, 2005), magma mixing (e.g.,Chamberlain et al., 2014;Lynn et al., 2017a), magma residence time before eruption (e.g.,Lynn et al., 2018;Nakamura, 1995), and subsolidus cooling rates of magmatic
bod-ies (e.g.,Coogan et al., 2007, 2002;Faak et al., 2013;Faak and Gillis,
2016;Sun and Lissenberg, 2018). Many theoretical and experimental studies have investigated the crystallographic and chemical parameters
controlling diffusion of major, minor and trace elements in mafic
min-erals, most particularly olivine, thus significantly improving the
avail-able diffusion coefficient database (e.g.,Coogan et al., 2005;Dohmen
and Chakraborty, 2007;Petry et al., 2004;Spandler and O'Neill, 2010). Concurrently, progresses in analytical techniques now allow measure-ment in situ of a broad range of chemical elemeasure-ments down to very low concentrations (<<ppm) such as Rare Earth Elements (REE) in olivine
(e.g.,Basch et al., 2018; D'Errico et al., 2016;Drouin et al., 2009;
Ferrando et al., 2018;Rampone et al., 2016;Sanfilippo et al., 2014).
These developments allow the collection of chemical profiles for many
elements that have different rates of diffusion, thus permitting detailed investigations on the timescales of complex processes.
In this study, we used geochemical profiles (Ferrando et al., 2018
and this study) in olivine because it is the only phase present from the mantle protolith to the reactive formation and cooling of olivine-rich
troctolites. Geochemical profiles record two processes: (i) flat profiles
of olivine forsterite content (Fo) and Ni, Mn, Co and Zn compositions indicate complete re-equilibration of olivine with the percolating melt
during melt-rock interactions; and (ii) convex profiles of Ca, Y and Yb
document subsolidus re-equilibration during cooling. However,
time-scales of the high-temperature (hereafter referred to as‘magmatic
temperatures’) reactive process and cooling remain to be estimated.
(i) We reproduce the measured homogeneous element profiles by
numerically modelling their chemical diffusive equilibration and re-trieve minimum timescale estimates of high temperature (> 1200 °C) re-equilibration. To account for diffusion anisotropy in olivine (D[001]> D[010]≈ D[100]for Fe\\Mg, Ni and Mn, D[001]>> D[100]> D [010]for Co, while Ca, Y and Lu are nearly isotropic;Table 1and
refer-ences therein), and to consider the influence of spatial dimensions,
crys-tal morphology, and sectioning, we used a three-dimensional (3D)
numerical model (Jollands and Müntener, 2019;Lynn et al., 2017b;
Shea et al., 2015). Simultaneous diffusive re-equilibration of eight
dif-fusing species comprising major, minor and trace elements (Fe\\Mg,
Ni, Mn, Ca, Co, Zn, Y, Yb) in olivine were modelled. Our results provide
thefirst constraints on the timing of melt-rock interactions largely
de-scribed at oceanic spreading centers.
(ii) Subsolidus cooling rates have been previously investigated in the oceanic crust at the Atlantis Massif OCC, although most constraints
doc-ument the latest stages of cooling (below 780°;Grimes et al., 2011;
Schoolmeesters et al., 2012) and cooling rates in the temperature
inter-val from magmatic conditions (~1200 °C) down to 900‐800 °C remain
unconstrained. In this study, calcium zoning in olivine is modelled
using the Ca-in-olivine geospeedometer (Coogan et al., 2007, 2002) to
determine ~1200 °C to ~1000 °C subsolidus cooling rates of Hole U1309D olivine-rich troctolites.
The results of this study characterize the timing of mantle incorpora-tion into the oceanic crust, which are used to reconstruct the thermal history and subsolidus cooling of the gabbroic sequence formed at an OCC, from high magmatic temperatures down to ~250 °C.
2. Geological setting and sampling
2.1. The oceanic crustal sequence at Atlantis Massif
Atlantis Massif is a domal structure located at 30°N on the western flank of the slow-spreading MAR, at the intersection with the Atlantis
transform fault (Fig. 1a). The corrugated core of the dome is a ~ 2 Ma
old OCC where oceanic crust is exposed via a long-lived and low-angle
detachment fault (e.g.,Blackman et al., 2002;Ildefonse et al., 2007).
On the central dome of Atlantis Massif, IODP Hole U1309D (Fig. 1a)
penetrated 1415.5 m below seafloor (mbsf) through a complex gabbroic
sequence (IODP Expedition 304/305;Blackman et al., 2011). The
recov-ered section is highly heterogeneous and comprises ~85% of gabbro, gabbronorite and olivine gabbro, with lesser oxide gabbro (7%) and minor MORB-type diabase intrusions (3%). The downhole
lithostratigra-phy is characterized by numerous inter-fingered intrusive bodies
vary-ing in thickness, with more evolved lithologies generally intrusive into
less evolved lithologies (e.g.,Blackman et al., 2006). Bulk Mg-values
Fig. 1. (a) Location of the Atlantis Massif OCC at 30°N on the MAR and topography of the Atlantis Transform Fault region (modified fromBlackman et al., 2006). (b) Examples of recovered intervals of (left) Ol-T1 (Core 248R-3, 82–104 cm) composed of wehrlitic (WEHRL), troctolitic (TROCT) and minor dunitic (DUN) domains, and (right) Ol-T2 (Core 248R-2, 5–26 cm) mainly composed of plagioclase-dunitic domains and cut by a gabbroic vein. (c) Downhole composition of IODP Hole U1309D (from left to right): 20 m running average of rock type recovered (white indicates no recovery), variations in whole-rock Mg # (Mg# (cationic ratio) = 100 × Mg/(Mg + Fetotal); modified afterBlackman et al., 2006), and Zircon Pb/U Ages where numbers indicate averages of measurements (Grimes et al., 2008). (d) Ni and Li average composition of olivine from single samples of Ol-Ts selected for this study in the interval 1100–1200 mbsf.
90 mol%, with local exceptions of lower values characterizing the most
evolved lithologies (Fig. 1c). No systematic compositional trends are
ob-served downhole (i.e., trend of fractional crystallization;Godard et al.,
2009). Such complex structural and textural relationships, together
with chemical evolution confined in discrete lenses and the variable
downhole zircon ages indicate that melt injections occurred at random
depths during the construction of the gabbroic sequence (Blackman
et al., 2006;Godard et al., 2009;Grimes et al., 2008). The total duration of magmatic accretion was inferred to be ~200 ka by radiometric dating
of zircons (Grimes et al., 2008;Fig. 1c) and occurred over two main
pe-riods, with an older intrusive event forming the deepest interval from 600 mbsf to the bottom of the hole (~1.24 Ma average) and a younger
event occurring throughout the core (~1.17 Ma average).Grimes et al.
(2008)inferred that emplacement of the oldest and deepest 635 m-thick lower crustal interval in Hole U1309D occurred within ~150 kyrs at a continuous growth rate of 1.6 cm/yr. These estimations, together with the recognition of over 250 intrusive igneous contacts throughout the Hole, allowed assessing an average thickness of single magma sills of
10 m and their emplacement every 630 yrs (Grimes et al., 2008).
Minor discrete intervals of serpentinized and locally impregnated
residual harzburgite (<1% of total recovery in Hole U1309D;Godard
et al., 2009; Tamura et al., 2008) were identified within the first
~200 m of Site U1309. Mantle rocks disappear downhole and ultramafic
intervals are represented by olivine-rich troctolites (Ol-T; 70–90%
olivine, 5–25% plagioclase, 5–25% clinopyroxene;Fig. 1b) occurring in
22 discrete intervals and covering ~5% of total recovery (Blackman
et al., 2006).
2.2. Hole U1309D olivine-rich troctolites
Ol-T is the dominant lithology in the interval between 1100 and
1300 mbsf (Fig. 1) and is locally very fresh with <1% serpentinization.
Drouin et al., 2009, 2010andSuhr et al. (2008)interpreted Ol-Ts from the Atlantis Massif OCC as resulting from melt-rock interaction in a
reac-tive porousflow process.Ferrando et al. (2018)demonstrated that the
multi-stage reactive process is triggered by melt infiltration into mantle
harzburgite. Reactive percolation of MORB-type melts leads to mantle
olivine dissolution, which locally modifies the composition of melts. In
turn, the tholeiitic crystallization suite is modified: concomitant
crystal-lization of interstitial clinopyroxene and plagioclase occurs whilst
oliv-ine is re-equilibrating with the migrating melt (Drouin et al., 2009;
Ferrando et al., 2018).Ferrando et al. (2018)posit that, as temperature
decreases, locally modified melts partially crystallize in cross-cutting
gabbroic veins andfinally form part of the gabbroic intrusions building
the crustal sequence at the Atlantis Massif OCC.
At the grain scale, during the open-system process of mantle-melt interaction, olivine is eroded (i.e., partial dissolution) and partially
pre-cipitated at crystal rims (Ferrando et al., 2018). The large variations and
non-systematic downhole correlations of Ni and Li contents (Ni =
1800–2820 ppm; Li = 1.5–2.9 ppm) in olivine from Ol-Ts (Fig. 1d) are
locally inherited from the precursor heterogeneous mantle (Fig. 10 in
Ferrando et al., 2018).Ferrando et al. (2018)defined two endmember Ol-Ts on the basis of their structural and textural characteristics, and
of their olivine composition. Ol-T1 (<77 vol% modal olivine;Table 2)
is the most reacted end-member (i.e. more mantle olivine dissolution) characterized by olivine grains showing Mg# ~ 85 mol% and Ni contents between 1870 and 2820 ppm, embedded in large oikocrysts of
plagio-clase and clinopyroxene (Fig. 4e in Ferrando et al., 2018). Ol-T2
(>77 vol% modal olivine;Table 2) is the least reacted end-member
with olivine displaying more evolved composition compared with
Ol-T1 (Mg# ~ 84 mol% and Ni = 1790–2130 ppm) and forming aggregates
with interstitial plagioclase and minor clinopyroxene (Fig. 4f in
Ferrando et al., 2018). In both Ol-T1 and Ol-T2 olivine compositions are in equilibrium with adjacent clinopyroxene and plagioclase
(Figs. 8–9 inFerrando et al., 2018).
Geochemical profiles in olivine (Fig. 2, Supplementary Figs. S1a and
S1b) were performed byFerrando et al. (2018)and in this study along
preferred crystallographic directions that were determined by Electron
Backscatter Diffraction analyses (EBSD). These profiles record two
dis-tinct processes that involved chemical diffusive re-equilibration,first
at magmatic temperatures (~1250 °C) during melt-rock interactions and subsequently at decreasing temperature under subsolidus condi-tions (down to ~1000 °C) during cooling.
At magmatic temperatures, the fast diffusive transport of elements
in olivine (e.g.,Dohmen and Chakraborty, 2007), which overall shows
less than three orders of magnitude difference in diffusivity (Table 1),
is able to reset olivine compositions. Consistently, Fo [Fo = 100 × Mg/
(Mg + Fe)], Ni, Mn, Co and Zn showflat profiles in all principal
crystal-lographic directions ([100], [010] and [001]; Supplementary Figs. S1a
and S1b). Flat profiles in olivine are interpreted to be the result of
complete diffusive re-equilibration with the reacted and modified
percolating melt during melt-rock interactions. These profiles can be
used to quantify the minimum timescales over which chemical re-equilibration likely occurred. This process of olivine re-re-equilibration is
herein referred to as‘diffusive re-equilibration at magmatic
tempera-tures (MT diffusive re-equilibration)’.
Although Ca in olivine was likely also completely re-equilibrated
during melt-rock interactions, Ca profiles show lower Ca contents at
the olivine crystal rim compared to the relative crystal core (Fig. 2).
Similar convex profiles are documented in some olivine crystals also
for Y and Heavy- REE (HREE, e.g., Yb; Fig. 2and Supplementary
Figs. S1a and S1b). Being preferentially hosted in clinopyroxene, Ca, Y and HREE can diffuse from olivine into clinopyroxene at subsolidus con-ditions due to the dependence of element partitioning on temperature
decrease (e.g.,Coogan et al., 2002;Witt-Eickschen and O'Neill, 2005).
In turn, the rate of diffusion decreases at decreasing temperature
(e.g.Coogan et al., 2005) hampering complete re-equilibration of olivine
core with the relative olivine rim at subsolidus conditions. We thus infer
that convex profiles do not result from the previous complete MT
diffu-sive re-equilibration during melt-rock interactions (i.e., leading toflat
profiles). They rather record subsolidus cooling after magma
emplace-ment (see discussion below). 3. Methodology
3.1. Geochemical profiles in olivine
The mineral modal contents, textures (grain size and phase
relation-ships) and microstructures analysed by EBSD are described inFerrando
et al. (2018). We focused this study on core-rim geochemical variations in 42 olivine grains from samples of Ol-T1 and Ol-T2, and we report complementary analyses of in situ major and trace element
composi-tions (Fig. 2and Supplementary Material Table S1). A subset of 12
most representative rim-core-rim profiles is shown in Supplementary
Figs. S1a and S1b. Absolute concentrations of Yb are the highest of the Rare Earth Elements (REE) in olivine (Supplementary Material
Table S1), therefore we report Yb inFig. 2as representative of olivine
REE concentrations.
Because element diffusion is overall anisotropic, geochemical pro-files were collected along preferred directions selected parallel to (at least) one of the three principal crystallographic axis of olivine having
axes plunge (measured from the sample surface) <5° (Ferrando et al.,
2018). All three crystallographic directions of olivine were investigated
in each sample. Core-rim and rim-to-rim traverses were measured with
19–80 μm spacing for major and minor elements, and 70–500 μm for
trace elements (Supplementary Material Table S1).
Major elements and trace elements were determined using analyti-cal instruments from the Microsonde Sud (Géosciences Montpellier, University of Montpellier). Major elements were measured by Electron Probe Micro Analyser (EPMA) using a CAMECA SX100 equipped with five wavelength-dispersive X-ray spectrometers (WDS); accelerating
Fig. 2. tChemical profiles across olivine crystals for Mg#, NiO wt%, Co (ppm), Zn (ppm), CaO wt%, Y (ppm) and Yb (ppm). Sample name and location of relative profile on single olivine are reported on top. Orange dots are used for Ol-T1 and red for Ol-T2. (For interpretation of the references to colour in thisfigure legend, the reader is referred to the web version of this article.)
potential was set to 20 kV, beam current at 10 nA, and counting times were 30 s for all elements. Natural minerals and synthetic oxides are used as standards. In situ trace elements were measured after EPMA analyses to minimize loss of sample. Surface cleaning was performed
on the selected EPMA spot by pre-ablation in 5–6 pulse per second.
We used a Thermo Scientific Element 2 XR (eXtended Range) high
res-olution - Inductively Coupled Plasma Mass Spectrometry (ICPMS). The ICP-MS is coupled with laser ablation (LA) system, a Microlas (Geolas Q+) automated platform with a 193 nm Excimer Compex 102 laser
from LambdaPhysik. An in-house modified 30 cm3ablation cell with a
helium atmosphere was used to enhance sensitivity and reduce
inter-element fractionation (Günther and Heinrich, 1999). The laser energy
density was set to 12–15 J cm2and repetition rate at 8–10 Hz. The
laser spot size was 102–77 μm. Data were collected in time resolved
ac-quisition mode with the background signal collected for 2 min followed by 1 min of sample ablation, or with 2.33 min for the blank and 40 s sampling. Concentrations were calibrated against the NIST 612 rhyolitic
glass using the values given inPearce et al. (1997). Data were reduced
with the GLITTER software package (Van Achterberg et al., 2001)
using the linearfit to ratio method, and29Si was used for internal
stan-dardization relative to EPMA data. Signals were carefully monitored and
data werefiltered for spikes on an element by element basis. Detection
limits were between 0.07 and 1 ppm for Ni, <55 ppb for Mn, Cu, Zn and <22 ppb for Co; they were <5 ppb for most incompatible elements (i.e., Y and HREE). Reference basalt BIR-1G was used as internal stan-dard to monitor accuracy as well as reproducibility within single series and between runs. This resulted in reproducibility better than 5% for Co
and REE, and it is <12% for all other elements (seeFerrando et al., 2018
for details and Tables).
3.2. Numerical diffusion modelling
Diffusion in solids is described using Fick's law (Ji= − Di∂Ci/∂x),
which states that the concentration evolution of a chemical species i
with time (flux Ji) is governed by its rate of diffusion (diffusion coef
fi-cient, Di, m2/s) and depends on the concentration gradient of i along a
given direction x (∂Ci/∂x), where Ciis the concentration of i. Most
nu-merical diffusion models are solutions of the Fick's second law (Crank,
1975). Several solutions of this equation have been published in the
lit-erature to evaluate the timescales of diffusion in minerals in various
al-though simplified magmatic systems.
Flat profiles of Fo, Ni, Mn, Co and Zn were modelled to quantify the
minimum timescales of MT diffusive re-equilibration (Fig. 2,
Supple-mentary Figs. S1a and S1b). To account for diffusion anisotropy in
oliv-ine, and to consider the influence of crystal size, morphology, and
sectioning, MT diffusive re-equilibration was simulated using
three-dimensional (3D) diffusion models modified afterShea et al. (2015).
Additionally, cooling rates were estimated using one-dimensional
models (Ca-in-olivine geospeedometry;Dodson, 1973;Coogan et al.,
2002, 2007) tofit CaO convex profiles measured in olivine (Fig. 2, Sup-plementary Figs. S1a and S1b). The choice of using a 1D model for Ca dif-fusion at subsolidus conditions is supported by the negligible dependence of Ca diffusion rates on the crystallographic orientation of
olivine (Coogan et al., 2005); the CaO profiles are convex in most olivine
crystals and in all measured directions, thus indicating that crystal shape and diffusion anisotropy (3D effects) did not affect Ca diffusion
at subsolidus conditions. Among CaO convex profiles, those displaying
flat plateaus at the crystal core were preferentially selected for models of Ca-in-olivine geospeedometry in this study. We avoided asymmetric
CaO convex profiles and those showing dipping plateau, as they are
common in sections that are off centre and/or oblique to principal crys-tallographic axes and can contribute to the uncertainty in retrieved
cooling rates (e.g.,Shea et al., 2015).
Data in this study are from olivines that span the range of observed grain sizes from 1 to 4 mm. In the following sections we illustrate the details of MT diffusive re-equilibration models and Ca-in-olivine geospeedometry.
3.2.1. MT diffusive re-equilibration models
Chemical re-equilibration in olivine was simulated for eight ele-ments that were chosen on the basis of their compatibility in olivine,
from compatible (Fe\\Mg [Fo], Ni) to moderately incompatible (Mn,
Co, Zn, Ca) and highly incompatible elements (Y, Lu). Trace elements were investigated as common tracers of melt-rock interaction
pro-cesses, during which their concentrations are strongly modified
(e.g.,Basch et al., 2018;Ferrando et al., 2018;Rampone et al., 2016;
Sanfilippo et al., 2014). Moreover, the modelled elements cover a wide
range of diffusion rates in olivine (Table 1) allowing characterization
of a broad range of minimum re-equilibration timescales.
Simulations were performed using three solutions of the 3D form of
Fick's second law (Crank, 1975). Solutions for anisotropic diffusing
species were applied for Fe\\Mg, Ni, and Mn using the
concentration-dependent equation (Eq.(A1)), and Ca using the non-concentration
dependent form (Eq.(A2)); trace elements were modelled using the
non-concentration dependent form for isotropic diffusion (Eq.(A3))
(see Appendix A for details on the equations). Because anisotropic diffu-sion of Co was documented exclusively at a single temperature
condi-tion (i.e., 1300 °C;Spandler and O'Neill, 2010) and the Arrhenius
relationship for DCois only available for Co diffusion along the
crystallo-graphic axis [001] (DCo[001], i.e., fast direction;Ito et al., 1999), we
modelled isotropic diffusion of Co using DCo[001]in Eq.(A3). We are
aware that this simplified approach possibly underestimates the
time-scales computed from Co profiles (i.e., use of the fastest DCo), but we
emphasize that the overall DCois comparable to the Diof the other
trace elements (seeTable 1) thus leading to timescales in the same
order of magnitude as Y and REE.
Although simple crystal geometries and analytical solutions are
often used to calculate timescales of element diffusion,Shea et al.
(2015)demonstrate that using spherical crystals lead to systematic
overestimations of timescales. Following thesefindings, olivine is here
modelled with polyhedral morphology (Shea et al., 2015) allowing to
reproduce a realistic olivine shape (Fig. 3). The best preserved olivine
crystals in the studied Ol-Ts have dimensions typical of euhedral olivine
crystals with c-axis >> b-axis > a-axis and c-axis≈ 2 * a-axis
(Supple-mentary Fig. S2). Comparable olivine crystal shapes have been previ-ously documented by Welsch et al. (2014) in samples from the same Hole U1309D. Therefore, we assumed a euhedral crystal similar to
Fig. 3. (a) 3D numerical olivine used in the models. Dashed white line indicates 2D section taken perpendicular to the c-axis [001]. Models were run twice, with 1 and 3 mm c-axis lengths. (b) Initial olivine (C0) and melt (Ci= C1) compositions used in the models. An
initially homogeneous crystal with a sharp compositional boundary to the surrounding melt was then allowed to diffuse at the conditions specified for each model and element.
olivines often described in erupted lavas (e.g.,Lynn et al., 2017a, 2017b). The 3D olivine + melt models have dimensions of 221x221x221 voxels (a voxel being a pixel in three dimensions). The modelled olivine had di-mensions of 201 voxels along the c-axis, 121 voxels along the b-axis,
and 95 voxels along the a-axis (e.g.,Shea et al., 2015;Fig. 3). Two crystal
sizes reflecting the minimum and maximum olivine sizes observed in
the thin sections were modelled with different voxel resolutions. For
small grains c-axis was set ~1 mm long (resolution of 4x4x4μm per
voxel; olivine crystal of dimensions x = 804μm, y = 484 μm, z =
380μm) and for coarser grains ~4 mm long (resolution of 20x20x20
μm per voxel; olivine crystal of dimensions x = 4020 μm, y =
2420μm, z = 1900 μm).
The advantage of the 3D model approach is that mineral core com-positions can be tracked through time. This provides insights into how resilient original mantle olivine compositions are to being overprinted
by secondary processes such as melt-rock interactions. A simplified
es-timate of diffusion distance using x = ffiffiffiffiffiffiffiffiffiffiðDitÞ p
has been applied by nu-merous previous studies to infer the time (t) required to generate a
diffusion profile of a given length (x) for an element of interest with
known Di. However, crystals are 3D objects subject to diffusivefluxes
in all dimensions. While this simplified 1D equation might broadly
char-acterize the time required for diffusive re-equilibration to reach a crystal's core, our 3D models allow us to determine how long it will take to completely re-equilibrate that core composition. Thus, this modelling approach is essential to understand how quickly minerals re-equilibrate with their surroundings and cannot be estimated using
the simplified 1D equation.
The continuous change of melt composition at the interface between crystal and melt during dissolution-precipitation processes would be best modelled using moving compositional boundary conditions
(e.g.,Chakraborty, 2008) combined with changes in crystal shape
(Chakraborty, 2018). However, experimental investigations of reactive percolation of a primitive melt (MORB-type melt) through a dunitic (Borghini et al., 2018) or troctolitic matrix (e.g.,Yang et al., 2019) docu-mented that at magmatic conditions (T ~ 1250 °C; similar to U1309D Ol-Ts formations) dissolution-reprecipitation reactions of olivine are rapid,
whileLiang (2003)in an experimental and theoretical study of the
ki-netics of melt-rock reaction demonstrated that diffusion of multiple
el-ements in basaltic melts (of the order ~10−11m2/s; e.g.,Kress and
Ghiorso, 1995; Liang, 2003) is faster than diffusive re-equilibration in
minerals (seeTable 1). Also, we have no constraints on (i) the variation
of melt volume due to dissolution-precipitation, (ii) the composition of olivine and melt at a given stage of the reactive process, and (iii) the ef-fect of changes in crystal shape on diffusion rates. Therefore, for simplic-ity, we modelled MT diffusive re-equilibration assuming static crystal shape and static boundary conditions (e.g., constant melt composition). This assumption implies instantaneous change in melt composition (i.e., fast element diffusion in melt; e.g., Liang, 2003) produced by melt-rock interactions and melt-olivine equilibrium at the crystal rim.
Diffusion simulations were run until complete re-equilibration with the boundary condition was achieved. 20 models were saved at equal intervals throughout the diffusion simulation to track the evolution of the crystal's composition with time. These models were then sectioned
through the olivine crystal's core perpendicular to the c-axis (Fig. 3).
The composition of the central pixel in the 2D section (Fig. 3) was
sam-pled at regular intervals throughout the model diffusion time to track
the core compositional evolution. We also sampled 1D numerical tra-verses parallel to the a-axis in the ideal 2D section to illustrate the
evo-lution of compositional profiles with time. The extent of MT diffusion is
expressed in terms of % re-equilibration (noted % req; seeLynn et al.,
2017b), which allows a direct comparison among diffusing chemical species with different absolute concentrations (e.g., ppm, wt%, mol%;
Fig. 4a,c):
%req ¼ðCinitial−CmeasuredÞ Cinitial−Cequilibrium
x 100 ð1Þ
where Cinitialis the composition of mantle olivine before the onset of
dif-fusion, Cmeasuredis the composition of the olivine core after a given time
(expressed in years, examples inFig. 4b), and Cequilibriumis the
composi-tion of olivine in chemical equilibrium with the surrounding melt. The magnitude of zoning in a 1D traverse parallel to the a-axis across
olivine diminishes as time elapses and the profile becomes difficult to
resolve analytically; distinctive inter-element zoning becomes less
likely and the crystal nears“complete re-equilibration” beyond the
ki-netic window. We introduce the“effective % req” to take into account
typical analytical uncertainties on each element (i.e.,σi= average
ana-lytical error of the modelled element i). The effective % req adjusts the
reported timescales with the analytical profiles that could realistically
be resolved in natural samples as they near 100% equilibrium with the boundary conditions. Using the effective % req, we consider complete re-equilibration when olivine core composition in the 3D model reaches Cequilibrium+σi(Table 3).
We ran the MT diffusive re-equilibration models at pressure of
2 kbar and oxygen fugacity at QFM (for Fe\\Mg, Ni, Mn, and Ca),
which are appropriate for the Atlantis Massif based on the estimated
depths at which the gabbroic sequence was formed (~7 km;Grimes
et al., 2008). Temperatures were set at 1230 °C for models of Fo, Ni,
Co, Mn, Ca (Fig. 4a), which correspond to the calculated crystallization
temperature of plagioclase (Drouin et al., 2009). For the listed major
and minor elements, models were also ran at 1210 °C and 1250 °C to
ac-count for uncertainties in the calculated timescales related to the +/−
23 °C uncertainty of temperature estimate fromDrouin et al. (2009).
From these models at three different temperatures we calculatedΔtT,
which is the uncertainty in modelled timescale related to the uncer-tainty in the temperature estimates. The unceruncer-tainty at the high end of
the T range (+ΔtT) is the timescale at 1210 °C– timescale at 1230 °C,
and the low end of the T range (−ΔtT) is the timescale at 1250 °C–
time-scale at 1230 °C. As no temperature-dependent equation is available for
calculating DZn, DLuand DY(single value at 1300 °C,Spandler and
O'Neill, 2010;Table B1in Appendix B), we also ran MT diffusive
re-equilibration models of all elements at 1300 °C and fO2= 10–8.3bars
(QFM-1; corresponding to the conditions of experiments inSpandler
and O'Neill, 2010;Fig. 4c).
3.2.1.1. Diffusion coefficients. For the MT diffusive re-equilibration
models, Diwere calculated at appropriate magmatic temperatures,
pressure and oxygen fugacity conditions using the Arrhenius
relation-ships reported inTable B1. Uncertainties in Diare related to
uncer-tainties in activation energies (~10%; e.g.,Costa et al., 2008), in turn
leading to uncertainties in computed timescales of diffusion. Di calcu-lated at temperatures higher or lower than the temperature at which
Fig. 4. Results of MT diffusive re-equilibration model to simulate timescales of melt-rock interactions at 1230 ± 20 °C (a-b) and 1300 °C (c) using C0a as starting olivine composition
(Table B2). (a-c) Evolution of the modelled % re-equilibration (% req) of olivine core as function of time of re-equilibration (years). Symbols represent olivine grain size of 1 mm and lines are for grain size of 4 mm. Colors distinguish each modelled element and reported in the legend. The time of re-equilibration is compared with the time of emplacement of a single 10 m thick sill (~630 years, light grey) afterGrimes et al. (2008). In (a) shaded areas in green, Co, and black, Ni, represent the ±Δ in time of re-equilibration (±ΔtT) at given %
req related to uncertainties in crystallization temperature (Drouin et al., 2009). They are reported for Co and Ni in 1 mm models only to provide examples; ±ΔtTof all other elements
can be found inTable 3. (b) 1D rim-to-rim profile evolution along the a-axis across 1 mm sized olivine (see section '3.2.1 MT diffusive re-equilibration models' for procedure of 1D profile sampling) are reported for Fo, NiO (wt%) and MnO (wt%) from MT diffusive re-equilibration model at 1230 °C.. Colored lines (elements are represented by colors as reported in the legend) highlight olivine composition after 50% of diffusive re-equilibration and olivine composition at effective % req (marked with * in the legend). For details on the MT diffusive re-equilibration models see the text and Appendix A. (For interpretation of the references to colour in thisfigure legend, the reader is referred to the web version of this article.)
Diwas experimentally constrained can have uncertainties of up to 3–4
orders of magnitude. In this study, Diwere calculated at temperatures
comprised in the temperature interval at which their Arrhenius rela-tionships were experimentally calibrated. This minimizes the
uncer-tainties in timescales, which are likely encompassed by ΔtT and
uncertainties for olivine initial composition (see later in this section). Also, greater uncertainties in timescales in these natural rocks are re-lated to the variable grain size of olivine crystals that are here treated modelling different crystal sizes.
To account for Fo variations, DFe-Mg(Fo)and DMnwere calculated
using the diffusion coefficient equations from Dohmen and
Chakraborty (2007), and DNifromPetry et al. (2004)(Table B1). As no
such concentration-dependent equations are available for the Ca and
Co, we calculated DCa and DCo using the Arrhenius relationships
(Coogan et al., 2005for Ca andIto et al., 1999for Co;Table B1) calibrated for olivine Fo ~90.
Diffusion of REE is currently matter of debate and vast discrepancies exist between trace element diffusivities in olivine with rates that span
over several orders of magnitude (10−19–10−20m2
/s byCherniak,
2010; ~10−15m2/s bySpandler and O'Neill, 2010). This range may be
re-lated to the different melt-powder compositions used in the
experi-ments and, specifically, on the activity of SiO2and Ti contents in the
melt (Supplementary Fig. S3; see discussions inJollands et al., 2016
andBurgess and Cooper, 2013). The source of diffusant used by
Spandler and O'Neill (2010)is a trace-element-doped MORB-type
melt containing 12 wt% of Al2O3and 5 wt% of TiO2, in contrast to the
synthesized mixture of REE (La, Dy, or Yb) aluminate and synthetic
forsterite powders used byCherniak (2010)that does not reproduce a
silicate melt (TiO2= 0 wt%). On the other hand, the experiments of
Spandler and O'Neill (2010)(i) also use a San Carlos olivine with
doped trace elements at levels that do not reflect natural olivine
simi-larly toCherniak (2010), and (ii) were conducted only at single
temper-ature, pressure and oxygen fugacity conditions, which prohibit the extrapolation to temperatures other than 1300 °C. Moreover, REE diffu-sion rates have been suggested to also depend on the mechanism of
dif-fusion and concentration of diffusing species in olivine (Chakraborty,
2018). Similar tofindings of Li diffusion (Dohmen et al., 2010), in
most natural systems REE may diffuse via two mechanisms that differ
in diffusion rate (Chakraborty, 2018); though, dependence of the active
diffusion mechanism on the concentration of the diffusing species has
never been defined, nor diffusion coefficients for these different
mech-anisms have ever been quantified.
Because only the experiments ofSpandler and O'Neill (2010)
in-clude the buffering of all chemical activities (e.g., aSiO2) and because
the computed melt composition in equilibrium with interstitial
clinopyroxene (seeFerrando et al., 2018) contains ~12–14 wt% of
Al2O3we posit that the DREE,Y(for DREEonly DLuis available) values
fromSpandler and O'Neill (2010)are the most appropriate for model-ling minimum timescales of re-equilibration after melt-rock reactions
at Atlantis Massif. For this reason, we selected DZn, DLuand DYsingle
values fromSpandler and O'Neill (2010). The use of fast diffusivities
may lead to underestimation of timescales, and our model results repre-sent minimum timescales of re-equilibration.
3.2.1.2. Initial and boundary conditions. Following the outcome of previ-ous studies (see Section 2.2) on the reactive origin of Ol-Ts from assim-ilated mantle, we assumed that the initial composition of olivine is that
of olivines in mantle harzburgites (Cinitial in Eq. (1)). Mantle
harzburgites sampled at modern ridges are compositionally heteroge-neous due to different extents of mantle melting and impregnation
pro-cesses (e.g.,D'Errico et al., 2016;Regelous et al., 2016;Tamura et al.,
2008). Also, among the little data available of in situ trace element
com-position of olivines in abyssal harzburgite (e.g.,D'Errico et al., 2016;
Regelous et al., 2016) none is from samples collected at the Atlantis Massif OCC. The Gakkel Ridge is a slow- to ultra-slow avolcanic spread-ing ridge, similar to the region of MAR where the Atlantis Massif OCC is
located. For these reasons, we explored the variation in timescales of MT diffusive re-equilibration using two different initial compositions (Table B2): (i) C0a is the average composition of olivine in harzburgites
from Gakkel Ridge (D'Errico et al., 2016;Regelous et al., 2016); (ii) C0b is
the composition of a single harzburgite olivine corresponding to that
used byFerrando et al. (2018)(values of Zn and Co are averages of
the most primitive olivines from the global database of abyssal
perido-tite compositions inRegelous et al., 2016). The resulting variation in
modelled timescale related to initial compositions is expressed as
Δtx=│t(C0a) - t(C0b)│, where t(C0a) and t(C0b) are timescales
calcu-lated for C0a and C0b, respectively.
Boundary conditions for the models are set to the measured olivine compositions, which represent the last reacted melt composition that is in equilibrium with adjacent clinopyroxene and plagioclase present in
Ol-T (Cequilibrium in Eq. (1)). Boundary melt compositions are
constrained by the averages of olivine composition in the U1309D Ol-T1 (Ferrando et al., 2018; summarized inTable B2), as representative of the most reacted Ol-Ts.
3.2.2. Ca-in-olivine cooling speedometry
After all melt has crystallized, the temperature dependence of
parti-tion coefficients (Kd) between minerals (e.g.,Coogan et al., 2002;
Witt-Eickschen and O'Neill, 2005) controls the re-distribution of elements during subsolidus cooling at the rims of adjacent phases, and a core-rim compositional gradient is produced within single crystals. The down-temperature dependence of Ca, Y and REE partitioning between olivine and clinopyroxene leads to the re-distribution of these elements from olivine rims to adjacent clinopyroxenes. Diffusion rates in olivine
also decrease with decreasing temperature (e.g.,Coogan et al., 2002;
Dohmen and Chakraborty, 2007), resulting in incomplete chemical re-equilibration of the crystal, in turn generating core-to-rim variations in slow-diffusing elements, such as Ca content. Calcium is often used
to retrieve timescales of subsolidus cooling (e.g.,Coogan et al., 2007;
Faak and Gillis, 2016). In this study, we determine the cooling rates and closure temperatures for 7 selected olivine grains that record a pro-gressive decrease of Ca, Y and REE contents toward the crystal rim,
ap-plying the Ca-in-olivine geospeedometry (Eq. (A4) in Appendix
A) following the methods ofCoogan et al. (2007). Here, the term
‘clo-sure temperature’ describes the low temperature at which no
apprecia-ble Ca diffusion occurs in the studied olivine (measured using the
geothermometer byKöhler and Brey, 1990; seeCoogan et al., 2007).
We used the DCafromCoogan et al. (2005)(Eq.(A5)in Appendix
A) as it is consistent with the most recent measurements byBloch
et al. (2019), who combined LA-ICP-MS, SIMS and atom probe (LEAP) analytical techniques. The Ca-in-olivine method assumes clinopyroxene
is an infinite reservoir for Ca, thus we selected only olivines in contact
with clinopyroxene.
Table 1
Selected olivine diffusion coefficients (m2
/s) reported in logarithmic scale; T≈ 1200–1300 °C and Fo ≈ 92. Diffusion coefficients were measured along the axis re-ported on top of each column ([100], [010], [001] are respectively a-axis, b-axis, c-axis). [1]Dohmen et al. (2007); [2]Petry et al. (2004); [3]Ito et al. (1999); [4]Spandler and O'Neill (2010); [5]Coogan et al. (2005); [6]Cherniak (2010); * single value at 1300 °C; ** average between [010] and [001]; *** anisotropy not considered.
Elements Olivine diffusion coefficients
[100] [010] [001] Fe-Mg [1] −17.10 −17.15 −16.32 Ni [2] −17.00 −17.02 −16.24 Co [3] −15.41 Co [4] −15.41 −15.35 −14.57 Zn [4] * −14.75 ** −14.75 ** Mn [2] −17.60 −17.57 −16.79 Ca [5] −17.31 −16.92 −16.74 Y [4] * −15.23 −15.22 −15.22 Lu [4] * −15.38 −15.32 −15.32
Because clinopyroxene has much higher Ca, Y and REE contents (here on average, CaO = 21 wt%) compared to those of olivine (here on average, CaO = 0.08 wt%), no core-to-rim variations are
generally observed in clinopyroxene (i.e., infinite reservoir
bound-ary condition).
Initial conditions for the diffusion model were homogeneous and set
to the Ca content measured in the core of the olivine crystal (Fig. 5and
Supplementary Fig. S4). Initial model temperatures were determined
using the partition coefficient for Ca between olivine and clinopyroxene
(Köhler and Brey, 1990) and the Ca content preserved in the olivine core plateau. This initial temperature was then used to calculate the initial
DCaused in the diffusion model (DCafrom Eq.(A5)in Appendix A).
After one model iteration, the time elapsed (∂t) was used to determine
the change in temperature (ΔT) as a function of an assumed cooling
rate. We recalculated the partition coefficient, temperature, and
diffu-sion coefficient for Ca after each time-step throughout the models to
ex-tract linear cooling rates; different cooling rates were imposed on the
models until a bestfit to the measured profile was obtained (Fig. 5
and Supplementary Fig. S4). The total time elapsed represents only the diffusion time recorded by the zoning generated between the initial and closure temperatures. Thus, they are minimum estimates of the du-ration of high-T cooling histories.
4. Results
4.1. Timescales of mantle assimilation into the crust
Timescales of diffusive re-equilibration after mantle-melt
interac-tions are reported inTable 3and inFig. 4at given % req. Because there
is no chemical zoning in olivine (except Ca;Fig. 2and Supplementary
Figs. S1a and S1b), we have no constraints on the time elapsed after complete re-equilibration (past the kinetic window) and therefore,
the determined timescales arefirst order minimum durations.
Time-scales can change as a function of temperature, initial conditions and grain size. The related variability of MT diffusive re-equilibration time-scales is reported in the following.
The 1230 °C 3D modelling of 4 mm olivine grains using C0a as
starting olivine composition show complete MT diffusive re-equilibration (i.e., effective ~100% req) within 220 years for all major
and minor elements (Fe\\Mg, Ni, Mn, Co, Ca;Table 3andFig. 4a).
Using C0b as starting composition for the same set temperature and
grain size provides timescales lower than 250 yr, leading toΔtX< 100 yr
(Table 3). Models ran at 1210 °C and 1250 °C allowed us to test the var-iability in timescales related to the uncertainty of the geothermometric
estimates (T = 1230 +/− 23 °C). The 1210 °C models yield the highest
values of minimum durations of diffusive re-equilibration among all
models. The variability in timescales (ΔtT) between models at 1230 °C
(taken as reference values) and at 1210 °C and 1250 °C ranges from
+/− 5 yr to +/‐40 yr between all investigated elements (Table 3). At
any given T, initial composition and % req, the time of re-equilibration of an element decreases with decreasing grain size (from 4 mm to
1 mm inFig. 4a,c), as equilibrium can be achieved more rapidly in
smaller crystals.
The models of MT diffusive re-equilibration at 1300 °C allow com-parison between all investigated elements including Zn, Y and Lu (see 3.1 MT diffusive re-equilibration models). They yield shorter timescales
compared to those simulated at 1230 °C (Fig. 4) due to the
T-dependence of Di(faster diffusivities at higher T). For olivine grains of
4 mm size and C0a as starting olivine composition, the 1300 °C 3D
models show that Zn displays the shortest re-equilibration time of
Table 2
List of samples and their principal textural characteristics and mineral chemical compositions. Ol = olivine, Spl = spinel, Cpx = clinopyroxene, Plg = plagioclase. Mg# = 100 × cationic (Mg/(Mg + Fe), with all Fe as Fe2+.
Sample IODP Depth (mbsf) Lithology Mineral modes (% EBSD indexed points) Mineral Chemistry Ol Plg Cpx Mg# Ol An Plg Mg# Cpx 305-U1309D-234R-1 W, 22–26 1124.97 Ol Troctolite 1 75 22 3 85.6 78.5 87.4 305-U1309D-241R-2 W, 89–91 1160.66 Ol Troctolite 2 84 13 3 84.4 74.7 86.6 305-U1309D-247R-3 W, 16–18 1190.05 Ol Troctolite 1 69 7 24 85.0 79.4 86.7 305-U1309D-247R-3 W, 76–81 1190.65 Ol Troctolite 2 81 13 6 84.6 74.4 85.8 305-U1309D-248R-2 W, 38–41 1193.52 Ol Troctolite 2 81 16 3 84.0 73.1 85.7 305-U1309D-248R-2 W, 43–48 1193.57 Ol Troctolite 2 80 15 5 84.1 76.1 86.1 305-U1309D-248R-3 W, 29–32 1194.78 Ol Troctolite 1 78 18 4 85.0 76.5 86.4 305-U1309D-248R-3 W, 36–38 1194.85 Ol Troctolite 1 64 6 30 85.1 74.7 86.3 305-U1309D-248R-3 W, 131–134 1195.80 Ol Troctolite 1 75 20 5 85.2 75.7 87.7 Table 3
Timescales (yr = years) of chemical re-equilibration, after mantle-melt interactions; % req is the effective % req (see text for details). Millimeters (mm) refer to the diameter of modelled olivine.ΔtTis the difference between timescales calculated at 1210 °C (t1210), 1230 °C (t1230), 1250 °C (t1250) (+ΔtT= t1210- t1230;−ΔtT= t1250- t1230).ΔtX(│ΔtX│) is the difference
between timescales calculated using two different starting olivine composition C0a and C0b.
Element 3D Model at 1230 +/− 20 °C 3D Model at 1300 °C Re-equilibration time (yr) - C0a Re-equilibration
time (yr) - C0b
Re-equilibration time (yr) - C0a
Re-equilibration time (yr) - C0b % req 1 mm (1230 °C) ΔtT (1 mm) 4 mm (1230 °C) ΔtT (4 mm) 1 mm 4 mm ΔtX (1 mm) ΔtX (4 mm) % req 1 mm 4 mm 1 mm 4 mm ΔtX (1 mm) ΔtX (4 mm) Fe-Mg (Fo) 98 25 +/− 5 190 +/− 5 30 225 5 35 98 18 120 16 200 2 80 Ni 97 27 +/− 7 100 +/− 20 21 150 6 50 98 13 84 10 140 3 56 Mn 99 26 +/− 7 220 +/− 40 32 230 6 10 99 20 126 18 215 2 89 Co 98 15 +/− 5 80 +/− 20 14 85 1 5 97 4.2 28 4.5 32 0.3 4 Zn 99 4.5 13.5 5 19.5 0.5 6 Ca 97 25 +/− 5 150 +/− 40 40 250 15 100 97 19 117 20 115 1 2 Lu 97 7.5 44 7 45 0.5 1 Y 99 6.4 39.5 6.7 40 0.3 0.5
~14 yr, followed by Co (28 yr), Y (40 yr), Lu (44 yr), Ni (84 yr), Ca
(117 yr), Fe\\Mg and Mn (120–125 yr) (Table 3andFig. 4c). The
con-struction of 2D sections (Fig. 3) also permits sampling of 1D
rim-core-rim profiles across the olivine models. With these data we can track
the evolution of the composition of olivine zoning patterns at regular in-tervals throughout the 1 mm 3D models at 1230 °C for representative major elements (Fo), and compatible (Ni) and moderately incompatible
(Mn) minor elements (Fig. 4b). Elements are re-equilibrated at
rela-tively constant rate up to 80% req (Fig. 4) of the olivine core. Fo content
decreases and Ni and Mn increase until 80% relative to the starting
com-position (C0a) within 15 years (Fig. 4b). To attain complete
equilibration other ~10 years are necessary, showing that the re-equilibration process progressively slows down as crystal nears com-plete re-equilibration (last ~20% req.), due to decreasing compositional
gradient (Fig. 4).
Regardless of the grain size, set temperature and initial composition,
all elements spanning from Fe\\Mg to incompatible REE (i. e., Lu) and Y
show complete re-equilibration in olivine within 300 years (Table 3and
Fig. 4). We recall that in MT diffusive re-equilibration models from this study we assumed a polyhedral morphology of olivine (i.e., euhedral crystal). On the other hand, models of crystals with more equant aspect ratios, and thus ultimately a smaller volume, would result in even
shorter re-equilibration timescales (e.g.,Shea et al., 2015).
Overall, the computed minimum durations of re-equilibration are
lower than the 630 yrs (light grey band inFig. 4) estimated for the
fre-quency of emplacement of 10 m-thick sills, and overall lower than the total ~150 kyrs emplacement of the oldest and deepest 635 m-thick lower crustal interval (from 600 to 1235 mbsf) in the Atlantis Massif
OCC (Grimes et al., 2008).
4.2. Timescales of subsolidus re-equilibration
Ca, Y and REE are preferentially hosted in clinopyroxene. Ca contents (in wt% from EPMA analyses) are systematically lower at olivine rim
compared to the crystal core (Fig. 2 and Supplementary Material
Table S1 and Supplementary Figs. S1a, S1b). Y and Yb concentrations (in ppm from LA-ICP-MS analyses) are either constant in single olivines, or slightly lower at the crystal rim in comparison with the
correspond-ing core (Fig. 2and Supplementary Material Table S1 and
Supplemen-tary Figs. S1a, S1b). Nonetheless, because of serpentinization along olivine edges, most geochemical analyses were performed at a
mini-mum distance of 50μm from olivine grain boundary leading to a gap
in data at the actual mineral/mineral interface. Moreover, the low con-centration of trace elements in olivine force setting large spot sizes
(102–77 μm in this study), which possibly prevent the analyses of
com-positional variations at the rim of the olivine crystal. As a consequence, we cannot preclude possible local chemical variations at olivine rims for
some elements, which can explain theflat Y and Yb profiles measured in
most discarded olivine grains and some selected olivines (Supplemen-tary Figs. S1a and S1b).
Core-to-rim chemical variations in olivine suggest that Ca, and to a lesser extent Y and Yb, diffused from olivine into the adjacent phases (i.e., plagioclase and clinopyroxene), as the effect of subsolidus diffusive
re-equilibration driven by a temperature sensitive Kdduring cooling of
the gabbroic sequence at the Atlantis Massif.
Partitioning of other elements, such as Mg\\Fe and Ni, is also
temperature-sensitive (e.g., Faak et al., 2013; Roeder and Emslie,
1970;Witt-Eickschen and O'Neill, 2005). Olivine acts as an infinite res-ervoir for compatible elements, nominally Ni and Mg, at decreasing temperature similarly to Ca in clinopyroxene. Cooling of gabbroic rocks can induce chemical re-distribution of Ni and Mg into olivine
from clinopyroxene and plagioclase, respectively (Faak et al., 2013;
Witt-Eickschen and O'Neill, 2005). If this was the case for olivines in Hole U1309D Ol-Ts, diffusion of Ni and Mg# from clinopyroxene and
plagioclase into olivine would have resulted in convex profiles of Mg#
in plagioclase (e.g.,Faak et al., 2013;Sun and Lissenberg, 2018) and Ni
in clinopyroxene during cooling. However, no zoning of Mg# or Ni con-tents is observed in clinopyroxene and plagioclase (Supplementary Fig. S5). Because element diffusion in clinopyroxene and plagioclase is
slower than diffusion in olivine (DCpx~ 10−19–10−21; e.g.,Van Orman
et al., 2001;Zhang et al., 2010; DMgPlag= ~10−16;Faak et al., 2013), Ni
and Mg# contents decrease at the very edge of plagioclase and clinopyroxene crystals without further concentration change (i.e., lack of element diffusion) toward the crystal cores. Such concentration
var-iations (likely at the order of fewμm) cannot be detected at the
resolu-tion of the laser spot size (77μm), and therefore it is plausible that
subsolidus re-equilibration of plagioclase and clinopyroxene was simply not detected in the studied Ol-Ts. On the olivine side, Mg# and Ni con-tents would increase of just a negligible amount because they are sub-stantially more concentrated in olivine than in the adjacent phases
(i.e., olivine infinite reservoir). At the resolution of EPMA and
LA-ICP-MS used in this study, these chemical variations are also not detectable. Moreover, olivine, clinopyroxene and plagioclase should record
equilib-rium temperatures close to 1000 °C if compositions were modified at
subsolidus conditions (e.g.,Coogan et al., 2002;Sun and Liang, 2014).
Mineral chemical compositions indicate that these phases in Hole U1309D Ol-Ts are in equilibrium at temperatures between 1190 °C
and 1205 °C with a melt showing Mg# = 58–65 mol% (Ferrando et al.,
2018). The high equilibrium temperatures, together with the
(appar-ent) lack of Ni and Mg zoning and olivine acting as infinite reservoir,
in-dicate thatflat profiles of elements preferentially hosted in olivine
record magmatic processes, and that those concentrations in the
stud-ied olivine were likely not modified, or increased to a negligible extent,
during cooling.
Initial temperatures recorded by olivine core Ca contents are on average ~ 1200 °C. Closure temperatures recorded by olivine rims range between 1014 °C and 1042 °C. To limit uncertainties on the com-puted closure temperatures, we also calculated equilibrium tempera-tures between olivine and clinopyroxene rims using the lattice strain
model for REE and Y distribution among mantle minerals (Sun and
Liang, 2014). These equilibrium temperatures range between 1000 °C and 1050 °C overlapping the computed closure temperatures. The
dura-tion required to generate a goodfit to the measured Ca profile
repre-sents a minimum timescale over which cooling might have occurred. The higher-T and lower-T (< 900 °C) history of cooling is outside the ki-netic window of Ca, and thus the total duration of cooling is longer than
that calculated here via Ca-in-olivine cooling geospeedometry (Fig. 5
and Supplementary Fig. S4).
Fig. 5. Ca profile in olivine from sample 305-U1309D-248R-3 W, 36–38. Plotted lines are Ca-cooling model at increasing time step as reported in the legend. The initial composition is represented by the dotted black line. The cooling rate calculate from this Ca profile corresponds to 0.004 °C/yr. For details on the Ca-in-olivine geospeedometry see the text.
The linear cooling rates obtained from the models span a range from
0.01 to 0.001 °C/yr (Fig. 6), with a mean value of 0.004 °C/yr (Fig. 5and
Supplementary Fig. S4).Fig. 6a displays data from this study compared
with downhole cooling rates of the oceanic crustal sequence from Hole
U1309D determined by combined U\\Pb zircon crystallization ages,
(U\\Th)/He zircon thermochronometry and multicomponent magnetic
remanence data (Grimes et al., 2011;Schoolmeesters et al., 2012). These
cooling rates were determined at lower temperatures, from 780 °C to
~250 °C. All data (this study;Grimes et al., 2011;Schoolmeesters et al.,
2012) overlap within error, showing that cooling rates are relatively
constant with depth and time over a large range of temperatures from
magmatic conditions to present (Fig. 6a).
5. Discussion
5.1. Preserved mantle chemical heterogeneities
Reactive processes through oceanic crustal gabbros modify the com-position of migrating melts and pre-existing crystal matrix
(e.g.,Boulanger et al., 2019;Lissenberg et al., 2013;Lissenberg and
Dick, 2008). The spatial extent of these chemical modifications depends on the intensity of melt-rock interactions, which are in turn strictly re-lated to the frequency of magma inputs and the melt/rock ratio
(e.g.,Basch et al., 2018;Higgie and Tommasi, 2012;Lambart et al.,
2019;Fig. 6). Continuous melt infiltration at high melt/rock ratios and reactive melt transport can lead to progressive over-enrichments in
in-compatible elements (Godard et al., 1995) from the deepest section of
lower oceanic crust to shallower levels (e.g., Hess Deep crustal
se-quence;Lissenberg et al., 2013). High melt/rock ratios enable chemical
homogenization and re-equilibration of the protolith minerals with
the reacted melt (Fig. 6b) at the scale of few tens of meters. Extensive
melt-rock interactions are able to shift the composition of percolating melts toward the composition expected for fractionation of MORB at
el-evated pressures (Lissenberg and Dick, 2008). On the other hand,
epi-sodic inputs of magma and low melt-rock ratios integrated over time lead to changes in chemical composition of percolating melt and crystal
matrix at local, centimeter-scale (Basch et al., 2018;Fig. 7c), having a
significant effect on the geochemical budget of erupted basalts
(Paquet et al., 2016).
The lack of up-hole over-enrichments (Fig. 1c) and the low-pressure
magmatic differentiation signature of Hole U1309D gabbroic section
suggest that percolating melts were overall not chemically modified
by reactive porousflow at Atlantis Massif. Assuming continuous melt
in-filtration (e.g., beneath fast-spreading ridges), fast MT diffusive re-equilibration timescales (i.e., at magmatic conditions) would be
respon-sible of chemical homogeneity of olivines over tens of meters (Fig. 7b).
In contrast, the composition of olivines is heterogeneous throughout
Hole U1309D, as evidenced by Ni and Li (Fig. 1d), suggesting that the
distribution offlow paths was confined into discrete intervals. The
strong variations in Ni and Li are inherited from pre-existing chemical
heterogeneity of the mantle protolith (Ferrando et al., 2018) and record
local re-equilibration with the modified melt after the reactive process
(Fig. 7c) at relatively high melt/rock ratios.
The presence of Ol-Ts at different depths throughout Hole U1309D (Fig. 1c), together with their inheritance of mantle heterogeneity and the fast MT diffusive re-equilibration of olivine crystals modelled in this study, point to episodic small-scale magmatism. Although esti-mated re-equilibration timescales are minimum durations, we docu-ment that Ol-Ts formed and cooled during the ongoing uplift of the Atlantis Massif OCC (see 6.3 Role of detachment faults on crustal cooling), suggesting that they remain shortly at depth after formation.
5.2. Depths of magma emplacement at Atlantis Massif
Multiple episodes of magmatic infiltration triggered partial
assimila-tion and incorporaassimila-tion of lithospheric mantle slivers into the oceanic
crust. Subsequently, the melts locally modified by mantle partial
assim-ilation migrated upwards and crystallized at variable depths to form
gabbroic intrusions. It is worth noting that modified melts likely
repre-sent only a portion of the parental magmas forming the crustal se-quence. Injections and crystallization of magma preserving their
primary composition (i.e., not modified by reactive processes) also
par-ticipated to the building of the lower oceanic crust at the Atlantis Massif OCC. In the following, we use geothermometric estimates and cooling rates recorded in rocks from Hole U1309D to reconstruct the history of formation and cooling of the oceanic crust at the Atlantis Massif OCC. Estimations of crystallization temperatures of the interstitial phases
in Ol-Ts from the Atlantis Massif (Drouin et al., 2009;Ferrando et al.,
2018) indicate that melt-rock interactions occurred at temperatures of
~1230 °C. At such magmatic temperatures, our MT diffusive
re-equilibration models (Table 3;Fig. 4) demonstrate that pre-existing
mantle olivine crystals are completely re-equilibrated with the reactive percolating melt in either the time-frame of a single melt input (i.e. <
630 yrs. assuming the emplacement of a 10 m-thick sill;Grimes et al.,
2008), or within the ~150 kyrs of emplacement of the 635 m-thick
deepest interval in Hole U1309D (i.e., between 600 and 1235 mbsf;
Fig. 6. Cooling rates (°C/year) obtained from Ca-in-olivine geospeedometry on Ol-Ts from this study (orange dots, Ol-T1, and red dots, Ol-T2) compared with cooling rates of slow-and fast-spreading oceanic crust. (a) Dots are downhole Hole U1309D cooling rates determined for gabbroic rocks by combined U\\Pb zircon crystallization ages and (U\\Th)/He zircon thermochronometry (Grimes et al., 2011) and multicomponent magnetic remanence data (Schoolmeesters et al., 2012). For the latter, we selected initial cooling rates (780 °C–250 °C) to ignore the effect of late-stage hydrothermal circulation, which buffer temperatures and decreases the cooling rates to ~0.0003 °C/yr (Schoolmeesters et al., 2012). The dotted blue box indicates the depth interval in (b). (b) Zoom-in of the studied interval (1190–1196 mbsf). Dotted black lines indicate constant downhole conductive cooling during uplift toward the surface modelled assuming temperatures of 0 °C at surface and 1300 °C at depth of magma emplacement (model fromCoogan et al., 2007). Models of conductive cooling are reported for uplift rates of 10 and 20 mm/year. (c) Ranges of cooling rates of the (i) slow-spread oceanic crustal section at Atlantis Bank OCC (black bars) retrieved using Ca-in-olivine geospeedometry (Atlantis Bank OCC and another OCC along the Mid Atlantic Ridge at 23°N (MARK area);Coogan et al., 2007) and thermochronometric data (John et al., 2004), and (ii) shallower section of the fast-spread crustal sequence at Hess Deep Rift (grey bar;Coogan et al., 2007 Faak and Gillis, 2016;Sun and Lissenberg, 2018). (For interpretation of the references to colour in thisfigure legend, the reader is referred to the web version of this article.)
